Stable isotope geochemistry of clay minerals

Abstract

The equilibrium H- and O-isotope fractionations can be approximated by the following equations which are based on experimental, empirical and/or theoretical data: Hydrogen: 1000 ln α kaolinite-water = −2.2 × 10 6 × T −2 − 7.7 Oxygen: 1000 ln α kaolinite-water = 2.76 × 10 6 × T −2 − 6.75 1000 ln α smectite-water = 2.55 × 10 6 × T −2 − 4.05 1000 ln α illite-water = 2.39 × 10 6 × T −2 − 3.76 The equilibrium H-isotope fractionation factors vs . 10 6 × T −2 for kaolinite and probably smectite and illite are monotonic curves between 350-0°C. More complex curves, with a minimum fractionation near 200°C, are probably influenced by surface effects and/or disequilibrium fractionations among the different hydrogen sites. The H-isotope fractionations between smectite-water increase by ~70‰ from Fe-poor montmorillonite to nontronite at low temperatures. The pore-interlayer water in smectite H-isotope fractionation at low temperatures is ~20±10‰. The presence of organic matter can modify both the δD value of the clay analysis and its ‘water’ content. Clays — kaolinite, illite, smectite and probably halloysite — tend to retain their D/H and 18 O/ 16 O ratios unless subjected to more extreme diagenetic or metamorphic conditions or special local processes. Kinetic information is still only qualitative: for comparable grain sizes, hydrogen exchanges more rapidly than oxygen in the absence of recrystallization. Low-temperature diffusion coefficients cannot be calculated with sufficient precision from the higher temperature exchange data. The H- and O-isotope studies of clays can provide useful information about their conditions of formation.